The influence of land cover change in the Asian monsoon region on present-day and mid-Holocene climate

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Influence of LCC in the Asian Monsoon region on climate

A. Dallmeyer and M. Claussen

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Biogeosciences Discuss., 8, 1693–1731, 2011 www.biogeosciences-discuss.net/8/1693/2011/ doi:10.5194/bgd-8-1693-2011

Biogeosciences Discussions

This discussion paper is/has been under review for the journal Biogeosciences (BG). Please refer to the corresponding final paper in BG if available.

The influence of land cover change in the

Asian monsoon region on present-day

and mid-Holocene climate

A. Dallmeyer1and M. Claussen1,2

1

Max Planck Institute for Meteorology, KlimaCampus Hamburg, Germany

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Meteorological Institute, University of Hamburg, KlimaCampus Hamburg, Germany

Received: 16 February 2011 – Accepted: 17 February 2011 – Published: 24 February 2011 Correspondence to: A. Dallmeyer (anne.dallmeyer@zmaw.de)

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Abstract

Using the general circulation model ECHAM5/JSBACH, we investigate the biogeophysical effect of large-scale afforestation and deforestation in the Asian mon-soon domain on present-day and mid-Holocene climate. We demonstrate that the applied land cover change does not only modify the local climate but also change

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the climate in North Africa and the Middle East via teleconnections. Deforestation in the Asian monsoon domain enhances the rainfall in North Africa. In parts of the Sa-hara summer precipitation is more than doubled. In contrast, afforestation strongly decreases summer rainfall in the Middle East and even leads to the cessation of the rainfall-activity in some parts of this region.

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Regarding the local climate, deforestation results in a reduction of precipitation and a cooler climate as grass mostly has a higher albedo than forests. However, in the core region of the Asian monsoon the decrease of evaporative cooling in the monsoon season overcompensates this signal and results in a net warming. A_fforestation has mainly the opposite effect, although the pattern of change is less clear. It leads to

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more precipitation in most parts of the Asian monsoon domain and a warmer climate except for the southern regions where a stronger evaporation decreases near-surface temperatures in the monsoon season.

When prescribing mid-Holocene insolation, the pattern of local precipitation change di_ffers. A_fforestation particularly increases monsoon rainfall in the region along the

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low River which was the settlement area of major prehistoric cultures. In this region, the effect of land cover change on precipitation is half as large as the orbitally-induced pre-cipitation change. Thus, our model results reveal that mid- to late-Holocene land cover change could strongly have contributed to the decreasing Asian monsoon precipitation during the Holocene known from reconstructions.

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1 Introduction

The effects of large scale deforestation and afforestation have often been analysed in climate model studies. In this context, the biogeophysical consequences of tropi-cal rainforest clearance in single regions like Amazonia, tropitropi-cal Africa or Southeast Asia as well as total tropical deforestation has widely been discussed (e.g.

Henderson-5

Sellers et al., 1993; Polcher and Laval, 1994; Gedney and Valdes, 2000; Werth and Avissar, 2002 and 2005a, b; Findell and Knutson, 2006; Hasler et al., 2009; Snyder, 2010). A few studies also deal with the impact of large-scale extratropical forest cover change on climate (e.g. Bonan et al., 1992; Chalita and Le Treut, 1994; Claussen et al., 2001; Bathiany et al., 2010). Most of these studies agree that tropical forest

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creases precipitation and cools the local climate compared to lower vegetation, mainly due to an enhanced evaporative cooling. Boreal forests have the opposite effect on temperature since they have a lower albedo than the often snow-covered ground and thus increase the absorption of energy at the surface.

How does climate react to large-scale forest cover changes in the Asian monsoon

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region? The Asian monsoon region comprises parts of the tropical as well as the temperate climate zone that can even be covered by snow during winter. Due to the heterogeneous orography, including high mountain ranges and parts of the Tibetan Plateau, potential land cover is very diverse. Rainforest grows next to temperate and montane forest; Alpine tundra is located there as well as temperate steppes.

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Monsoon circulations are primarily induced by the thermal contrast between the con-tinent and the ocean, generating large scale pressure gradients (Webster et al., 1998). Therefore, monsoon systems have to be seen as coupled interactive atmosphere-ocean-land phenomena (Yasunari, 2007) in which the land surface plays an equally important part as the ocean. Since land cover changes affect the energy and water

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The Asian monsoon system is the strongest monsoon system worldwide. Including large parts of China and India, nearly two thirds of the global population live in the region influenced by this monsoon (Clift and Plumb, 2008). Changes in the monsoonal climate are therefore a matter of particular social, economic and environmental interest. In the last decades the Asian monsoon region experienced large anthropogenic land

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cover changes. Huge parts of the natural forest in eastern China have been converted into farmland (e.g. Pongratz et al., 2008). Steppe areas are severely affected by over-grazing and over-farming. Originally vegetated areas in northern China have been overrun by the rapidly expanding deserts. The impact of these land cover change has been analysed in several climate model studies (e.g. Zheng et al., 2002; Fu, 2003;

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Gao et al., 2007; Zhang and Gao, 2008). They all agree that the vast degradation of the land cover has caused a regional climate change. However, amplitude, sign and distribution of precipitation and temperature change differ and depend on the type of model (regional or global climate model) and the methods for applying the land cover change.

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South and East Asia are the homelands of some of the oldest human civilizations. These cultures evolved rapidly and often had advanced agricultural technologies and a well organised urban infrastructure. Particularly the Indus Valley (today’s western Pakistan and eastern Afghanistan) and the region around the Yellow River (Fu, 2003; Clift and Plumb, 2008) experienced a long continuing settlement history, starting in the

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early-Holocene (approx. 7000 BC and 6000 BC, respectively). The linkage between the development or collapse of theses cultures and the monsoon climate variability as well as the role early human societies could have played for the Holocene climate and environmental change are of special interest in the context of recent global warming. Some studies relate the decline of diverse major prehistoric cultures to intense

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Pollen-based vegetation reconstructions suggest that large parts of Central and Eastern China were much more covered by forest during the mid-Holocene. For in-stance, Ren (2007) suggested an up to 92% increased forest cover in the middle and lower reaches of the Yellow River for 6k (k=1000 yr before present) compared to today, which strongly declined in the following four millennia. According to Ren (2000), the

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spatial evolution of forest decline in China correlates well with spreading of cultures out of this area. However, Ren (2000) also remarks that this relation can only be a preliminary conclusion due to the lack of well-dated and well-resolved pollen data.

The steppe-forest boundary in East Asia was shifted north-westward by up to 500 km at 6k (Yu et al., 2000). Thus, the area around the present-day East Asian monsoon

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margin experienced substantial decrease of vegetation in the course of the Holocene. In eastern Inner Mongolia, pollen assemblages reached the maximum Holocene tree pollen fraction (40–60%) between 8k and 5k . Tree pollen decreased in this region to less than 20% within 2000 yr. Also for the nowadays semi-arid north-western Loess Plateau tree pollen dominate the early- to mid-Holocene pollen assemblages, reaching

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80% between 8.5k to 6.5k. Afterwards, land cover gradually changed to desert with sparse steppe in this region (Zhao et al., 2009).

Less information is known about the vegetation change in the Indian monsoon mar-gin (e.g. Pakistan and north-western India), but also for this area a vegetation loss during the Holocene is supposed (Singh et al., 1990; Ansari and Vink, 2007; Ivory and

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L ´ezine, 2009).

These examples of vegetation reconstructions point to the fact that the vegetated area in the Asian Monsoon domain was not only characterised by more forests during mid-Holocene but also expanded further inland.

The primary aim of this study is to assess the influence of large-scale land-cover

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is then presented on the basis of precipitation and temperature differences to a con-trol run with present-day potential vegetation. Secondly, we analyse the remote e_ffect these land cover changes have on precipitation in North Africa and the Middle East. Thirdly, we address the question what information we can deduced from these results regarding the relation of mid-Holocene climate, vegetation and neolithic cultures in the

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Asian monsoon region.

2 Model and experiments

To analyse the influence of large-scale land cover change on the Asian monsoon cli-mate, different sensitivity experiments are performed. In this study, we use the general circulation model ECHAM5 (Roeckner et al., 2003) coupled to the land-surface-scheme

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JSBACH (Raddatz et al., 2007). Both models have been developed at the Max-Planck-Institute for Meteorology. ECHAM5 run with 31 vertical levels and a spectral resolution of T63, which corresponds to a grid-box width of 1.875◦_{. JSBACH di}

fferentiates eight plant-functional types (PFTs). Forests can contain tropical and/or extratropical trees, which are either evergreen or deciduous. Shrubs are distinguished as raingreen shrubs

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or cold shrubs. Grass is classified as either C3 or C4 grass. The land surface in JSBACH is tiled in mosaics, so that several PFTs can cover one grid cell. Each grid cell also contains non-vegetated tiles representing the fraction of seasonally bare soil and permanently bare ground (desert). The models have been tested against obser-vations and reanalysis data proving that they capture the major structure of global and

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regional climate (e.g. Cui et al. 2006).

In our model setup, sea surface temperatures and sea-ice cover are prescribed according to pre-industrial conditions. Both have been extracted from a coarse resolution experiment, performed with the comprehensive Earth System Model ECHAM5/JSBACH-MPIOM. This model includes the dynamic ocean model MPIOM

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Under these boundary conditions, three experiments with differently prescribed land cover in an extended Asian monsoon region (see Fig. 1) are conducted: These are a control run with standard, natural land cover (CTRL), an experiment with forest in the entire region (FOR) and an experiment with grass cover in the entire region (GRA). Thereby, the region was selected combining the following criteria: We extend the Asian

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monsoon domain by including areas where previous model results have suggested a strong increase of vegetation for mid-Holocene orbital conditions to get a represen-tative monsoon domain for the Holocene (cf. Dallmeyer et al., 2010). Vegetation is not changed in high-elevated areas where the mean orography in the model exceeds 4500 m. Furthermore, the region incorporates areas where humans settled in

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and/or mid-Holocene, i.e. the region includes the Indus Valley, Northeast China and the Chinese provinces Shandong, Hebei, Shanxi, Shaanxi and parts of Gansu and Qinghai.

In the a_fforestation experiment (FOR), land cover is prescribed as forest in the entire region. In those grid-boxes that have already been forested in the standard land surface

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map, the ratio between the different tree types is kept constant and the forest sum is scaled to 100% fraction per grid-box. In grid-boxes without forest vegetation in the standard map, forest is prescribed to 50% evergreen and 50% deciduous trees. Trees are prescribed as extratropical or tropical to the north or south of 30◦_{N, respectively.}

The same method is applied in the deforestation experiment (GRA). For grid-boxes

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already containing grass PFTs in the standard map, the ratio between C3 and C4 grass is fixed and total grass sum per grid-box is scaled to 100% coverage. Regions not containing grass PFTs in the standard map are covered by 20% C3 and 80% C4 grass south of 30◦

N and 80% C3 and 20% C4 grass north of 30◦

N.

The difference between the land cover in the afforestation or deforestation

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ment and the control run can be seen in Fig. 2 for the vegetation types forest, shrubs and grass.

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Each experiment spans 102 yr, the first two years are considered as spin-up time and excluded from our analyses. The presented plots only show results exceeding the 95%-significance level of a standard student’s T-Test.

3 Results

The period of a reversed upper-tropospheric temperature gradient between 5◦

N and

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30◦_{N is often taken as a definition for the Asian monsoon season (Li and Yanai, 1996;}

Webster et al., 1998; Ye and Wu, 1998). According to this definition, the length of the monsoon season in our experiments remains unaffected regardless of which land cover is prescribed (not shown). Based on the month of earliest simulated onset and latest simulated withdrawal in the Asian monsoon domain we therefore divide our analysis

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period into two di_fferent seasons. These are the monsoon season (lasting from May to October) and the dry/cold season (lasting from November to April). Furthermore, we concentrate on precipitation and near-surface air temperature as the main parameters describing monsoonal influenced climates. The results presented in this study are based on 100 yr-means, averaged over the respective season.

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3.1 E_ffect of land cover change on precipitation

3.1.1 Afforestation experiment

The significant effect of a prescribed forest cover on seasonal mean precipitation is illustrated in Fig. 3. The land cover change predominantly has a local and only weak e_ffect on precipitation. Monsoon precipitation is increased particularly on the

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south-central Tibetan Plateau (up to 0.77 mm day−1_). _{In north-eastern China}

(ap-prox. 0.35 mm day−1_{) as well as at the monsoon margin zone in Inner Mongolia}

(ap-prox. 0.17 mm day−1_{), precipitation is also enhanced. In contrast, the western}

Ti-betan Plateau receives less precipitation (up to 0.26 mm day−1_{) in the a}

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experiment. At the Yangtze delta and parts of the Great Plain of China as well as in South India rainfall is reduced by up to 0.35 mm day−1

and 0.43 mm day−1

, respectively. In the dry/cold season, the land cover change causes a significant increase of precipitation on the Tibetan Plateau (ca. 0.39 mm day−1_{) and in north-eastern China}

(ca. 0.09 mm day−1

). Less precipitation occurs in the region west and north of the Tian

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Shan (up to 0.23 mm day−1_).

3.1.2 Deforestation experiment

Complete grass cover in the Asian monsoon region instead of present-day natural vegetation leads to a decrease of monsoon precipitation in a broad band in South and East Asia, including the Yangtze-Huanghe-plain, Indochina, Bay of Bengal, parts

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of the Tibetan Plateau and East India (Fig. 3). The maximum reduction in monsoon rainfall (2.3 mm day−1

) occurs on the south-eastern edge of the Tibetan Plateau. In contrast, monsoonal precipitation increases slightly, but significantly on the western Tibetan Plateau, reaching 0.26 mm day−1_{at most.}

In the dry/cold season, the strong precipitation reduction is confined to Eastern China

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(up to 0.56 mm day−1_{) and the Indochinese Peninsula (up to 0.35 mm day}−1_{). The}

cen-tral and western Tibetan Plateau receive more precipitation, in particular the western Himalaya with up to 0.3 mm day−1_.

3.2 Effect of land cover change on near-surface air temperature

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Figure 4 shows the significant effect of afforestation on near-surface air temperature. In the monsoon season, the regions north of 30◦_{N that experience a strong a}

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to decrease. The land cover change leads to significant cooling in some parts of the Indian West coast and southeast of the Tibetan Plateau (up to₋0.3 K).

In the dry/cold season, 2 m-temperature is particularly increased in Northeast China (up to 2.2 K) and on the southern flank of the Tibetan Plateau (up to 2.1 K). Eastern China, including the eastern Tibetan Plateau, North India, and the region north of

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the Arabian Sea experience a warmer climate (up to 0.4 K). Only at a few grid-boxes southeast of the Tibetan Plateau temperatures slightly decrease by up to−0.25 K.

In the deforestation experiment, monsoon season temperatures are lower in Northeast China (up to 0.35 K), Inner Mongolia (ca. 0.2 K) and on the Tibetan Plateau (up to

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0.55 K) compared to the control run (Fig. 4). In contrast, the deforested regions south of 25◦_{N experience a warming. Temperatures increase by up to 0.9 K in northern}

Thailand. Higher temperatures are also found in eastern China, particularly at and around the Great Plain (up to 0.6 K).

With prescribed grass cover, dry/cold season temperatures are lower in whole

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eastern China. The cooling is particularly pronounced in regions, which experience a strong deforestation, e.g. the Great Plain (ca. 0.7 K). Nevertheless, the strongest tem-perature decline occurs in Northeast China (up to −2.2 K). Climate becomes slightly

warmer in the north of India by up to 0.25 K. At the deforested southern flank of the Tibetan Plateau near-surface air-temperature decreases (ca.−0.8 K).

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4 Discussion

The land surface can force the atmosphere via the physical quantities snow cover, soil moisture and vegetation (Yasunari, 2007). In the context of this study, the effect of vegetation change on climate is most important, though vegetation change can also indirectly influences the other parameters. The land cover controls the energy balance

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at the surface via albedo and roughness length and influences the evapotranspiration. Modifications of the energy fluxes can lead to changes in temperature, pressure field and circulation (Pielke et al., 1998). Modifications of the evaporation due to land cover changes together with circulation changes affect the precipitation distribution. There-fore, we start this chapter with a discussion of the energy and water balance change

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related to the applied afforestation and deforestation. 4.1 Change of circulation and water balance

The Asian monsoon circulation is primarily driven by the seasonal differential heating between the Asian continent and the Indian and Pacific Ocean. During the summer

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monsoon season, a deep surface heat low covers the region between the Arabian peninsula and the Chinese East coast with lowest pressure in North India and Pakistan. In consequence of the afforestation, the monsoon trough deepens by up to−120 Pa in

North India (Fig. 5c). Therefore, the Indian monsoon is slightly strengthened, which at least contributes to the enhanced precipitation over southern Himalaya.

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In North and Northeast China, the precipitation increase coincides with an enhanced cloudiness (up to 3%, Fig. 5a) and an increased evaporation (up to 0.56 mm day−1_{, see}

Fig. 5b), probably indicating an enhanced water recycling in that region. The shift in cloud cover from the Great Plain – Japan region to North/Northeast China suggests a shift of the East Asian summer monsoon rain band, i.e. a north-westward extension or

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shift of the region influenced by the monsoon. This would also explain the reduction of precipitation at the Great Plain and surroundings. The approximation for the moisture convergence (precipitation-evaporation (p-e), Fig. 5d) reveals a negative anomaly (i.e., a moisture source) in that region. The extension of the monsoon is consistent with the wind field showing a slight enhancement of the East Asian summer monsoon flow.

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During the dry/cold season, the high pressure system in North China is weak-ened by up to −100 Pa (Fig. 6c). Therefore, the strong southward flow in eastern

China characterising the East Asian winter monsoon is also weakened. The pressure anomaly induces an anticyclonic wind anomaly with a core in the Yellow Sea which brings warm and wet air-masses to Northeast China. This seems to lead to more

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clouds and more precipitation in that region.

The increase of precipitation on the Eastern Tibetan Plateau coincides with an en-hanced cloudiness (up to 4%) and an increased evaporation (up to 0.38 mm day−1

) as well as a mean sea level low pressure anomaly (up to−220 Pa). The latter leads to

a wind anomaly from the Bay of Bengal to the continent, probably contributing to the

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precipitation change.

Deforestation leads to an attenuation of the Indian and East Asian summer monsoon flows (Fig. 7c). This circulation change is associated with a high pressure anomaly forming in South and Central Asia, in particular on the Tibetan Plateau (up to 110 Pa).

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High pressure anomalies are often related to descending motion and thus indicate less favourable conditions for the formation of precipitation. As a result, monsoon related convection, precipitation and also evaporation in south/south-eastern Asia is reduced (Fig. 7). Total cloud cover decreases by up to 5%, evaporation decreases by up to 0.78 mm day−1_{. In contrast, the increase of precipitation on the western Tibetan} 20

Plateau is accompanied by more evaporation (up to 0.3 mm day−1_{) and more clouds}

(up to 4.5%).

During the dry/cold season the moisture flux is more divergent in eastern China than in the control experiment (Fig. 8). In addition, evaporation rate and cloud cover are reduced by up to 0.52 mm day−1 _{and up to 4%, respectively. Both indicate a weaker} 25

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4.2 Change of albedo and turbulent heat fluxes

In the monsoon season, afforestation leads to a decrease of albedo by up to 9.5% (Fig. 9a). This results in an increase of the absorbed incoming shortwave radiation, which explains the increase of near-surface air temperature in Northeast China, Inner

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Mongolia as well as in the region southwest of the Tibetan Plateau. The additional shortwave radiation at the surface is mainly transported to the atmosphere by the sen-sible heat flux. The region of decreased near-surface temperature southeast of the Tibetan Plateau is associated with a region of less net shortwave radiation at the sur-face (ca. 6.5 W m−2_{), which is probably induced by more clouds (not significant) and a} 10

slight albedo increase in that region. At least partly, the increased evaporation/latent-heat flux (up to 5 W m−2

) and the decrease of sensible heat flux (up to 8.5 W m−2

) contribute to lower temperatures in this region due to an increase in evaporative cool-ing (Fig. 9).

The strong warming in Northeast China and on the Tibetan Plateau in the dry/cold

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season is induced by a strong albedo decrease (up to 20% and 35%, respectively, Fig. 10a) which leads to a strong increase of net shortwave radiation at the surface (up to 18 W m−2_{and 55 W m}−2_{). These regions are covered by snow in winter, so that}

the snow-masking effect of the additional forests might contribute to the strong albedo decrease.

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In the region southeast of the Tibetan Plateau, the temperature decrease coincides with less net shortwave radiation at the surface (up to−6 W m−2), which probably

re-sults from an enhanced cloud cover (up to 3%). 4.2.2 Deforestation experiment

The replacement of forest with grass cover in the Asian monsoon domain increases the

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air temperatures in Southeast China and the regions south of 25◦_{N are higher in the}

deforestation than in the control experiment. This is mainly a consequence of less precipitation/evaporation and less cloud cover in that region (cf. Fig. 7a and b) which overcompensate the e_ffect of increased albedo on the near-surface temperature.

The lower temperatures on the Tibetan Plateau are induced by less net solar energy

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at the surface (up to−11 W m−2), mainly due to more clouds (up to 4.5%). Stronger

evaporative cooling contributes to the lower temperatures on the western Himalaya. In Northeast China, the temperature decrease results from both, less net shortwave radiation at the surface (−6 W m−2) due to an albedo increase of up to 5% as well as a

higher evaporative cooling.

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In the dry/cold season, deforestation leads to an up to 25% increase of albedo in the eastern part of Northeast China (Fig. 12), which results in less net shortwave ra-diation at the surface (up to −20 W m−2). This part of China is usually covered by

more snow than the other parts of the Asian monsoon domain. Therefore, the lack of snow-masking by trees as well as a possibly enhanced snow-albedo feedback in the

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deforestation experiment may play an important role in decreasing the near-surface-temperature in this region.

4.3 Comparison to other model studies

So far, no other publications exist prescribing complete forest or grass cover in the Asian monsoon domain. But, our results show similar climate signals received in

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vious studies performing land cover change experiments. Like in studies with tropical deforestation, precipitation in our simulation decreases when grass cover is prescribed instead of potential present-day vegetation (cf. e.g. Werth and Avissar, 2005a). These studies also report that less forest cover in tropical regions leads to a warmer local climate (cf. e.g. Polcher and Laval, 1994). Increased near-surface air temperatures

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are also simulated in our deforestation experiment in the regions with a rather trop-ical climate, i.e. the region ca. south of approx. 26◦

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forest loss. This response is in line with previous studies of extratropical deforestation (Bonan, 2008). The results from our a_fforestation experiment reveal a similar pattern of precipitation change than model studies which analyse the effect of present-day land use change in China. For instances, Gao et al. (2007) have compared results from simulations with potential vegetation cover and observed land cover in China by using

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a regional climate model. Their results show that the vegetation conversion from forest to farmland reduces summer precipitation in northern and western China and enhance precipitation in Southeast China. Similar results were deduced by Zheng et al. (2002) by prescribing a simultaneous vegetation degradation in the northern Chinese prairies and the southern evergreen forests. Furthermore, precipitation increases at the Indian

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East coast. The forest loss between our a_fforestation experiment and the control run leads to a similar precipitation change in the monsoon season.

4.4 Remote e_ffect on precipitation in North Africa and the Middle East

Rodwell and Hoskins (1996) showed in an idealised model study that Asian monsoon related diabatic heating can induce descent in south-western Asia, the Arabian

Penin-15

sula and northern Africa via the initiation of Rossby-waves. Other studies emphasise the important role of diabatic heating on the Tibetan Plateau by forming desert climate in North Africa, Central Asia and the Middle East (e.g. Ye and Wu, 1998; Duan and Wu, 2005). As a consequence of this link, land cover change in the Asian monsoon domain may affect the climate in these regions as it has a strong influence on the

en-20

ergy balance. We address the possible remote e_ffect of land cover change by analysing precipitation change in the monsoon season.

Figure 3 already reveals that the land cover change prescribed in this model study has predominantly a local and only weak influence on precipitation regarding the change of absolute precipitation amount. But, when looking at the relative precipitation

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around the Southern Caspian Sea. Precipitation is more than halved in a large area. In many places, precipitation even decreases by 80–100%. The upper-tropospheric velocity-potential anomaly suggests that this change results from an increase in upper-level wind convergence, enhancing the descending motion and suppressing convective precipitation. In contrast, no significant large-area precipitation change occurs in the

5

North African monsoon region and the Sahara.

Deforestation in the Asian monsoon region leads to a significant increase of precip-itation in large parts of North Africa. In the Sahel zone precipprecip-itation amount rises by more than 10%, in the eastern Sahel even by more than 50%. The west coastal regions receive more than 25% more precipitation. In the central Sahara precipitation is more

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than doubled. Besides, precipitation rate is increased in the north-eastern Mediter-ranean by 20%–30%. Figure 7 shows that the deforestation leads to a low-pressure anomaly of up to 33 Pa in Northern Africa. Therefore, the heat low in the Sahara is ex-panded and deepened. This induces an enhanced low-level flow from the ocean to the continent, bringing more moisture to the Sahara and Sahel. Furthermore, a reduced

15

convergence in the upper troposphere above Africa and the Mediterranean (Fig. 13) facilitate the formation of deep convection and precipitation.

4.5 Effect of idealised forest-cover change on mid-Holocene climate

During mid-Holocene, large parts of the Asian monsoon domain were covered by more forest and the forested and vegetated region expanded further inland compared to

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present-day (e.g. Yu et al., 2000; Ren and Beug, 2002; Ren, 2007; Herzschuh et al., 2010). Pollen-based vegetation reconstructions reveal broad land-cover degradation since that period (e.g. Zhao et al., 2009). These changes are often related to the decrease in Asian monsoon strength (e.g. Fleitmann, 2003; Wang et al., 2005; Maher, 2008) entailing shifts in precipitation pattern.

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could strongly have influenced the development and decline of different prehistoric cultures in the Asian monsoon domain (cf. Fu, 2003; Clift and Plumb, 2008). The linkage between mid-Holocene climate, vegetation and human dynamics in the Asian monsoon region is therefore a scientific question worth to investigate.

As the Asian topography is very complex, the relatively sparse-distributed

recon-5

structions can not necessarily represent the land cover in the entire region. Numeri-cal experiments, for example conducted with the comprehensive Earth system model ECHAM5/JSBACH-MPIOM support the increase of vegetation and forest in the Asian monsoon domain under 6k climate conditions (Dallmeyer et al., 2010). But, compared to the available pollen records, the magnitude of Holocene vegetation change in these

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simulations seems to be underestimated. A complete picture of the mid-Holocene land cover in the Asian monsoon region and its change during the following millennia thus cannot be derived yet, neither from reconstructions nor from previous model-results.

To assess the biogeophysical impact of Holocene land cover change on the Asian monsoon climate we therefore chose an idealised approach: We assume that in

mid-15

Holocene, the entire Asian monsoon domain was covered by forest and that mid- to late-Holocene climate change as well as early human forest clearance have altered the vegetation up to its present-day potential distribution. In other words, we repeat the afforestation and control experiment with same boundary conditions with the exception that we prescribed mid-Holocene (6k=6000 yr before present) insolation (FOR6k and

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CTRL6k). The biogeophysical effect of the idealised Holocene land cover change is then defined as difference between the experiments “FOR6k” and “CTRL6k”.

Figure 14a depicts the summer monsoon related precipitation difference between the afforestation and control experiment under 6k-orbital conditions. Overall, the am-plitude of monsoon precipitation change is similar as for 0k-orbit (cf. Fig. 3, upper

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less precipitation in south-eastern China (up to 0.2 mm day−1_{) and significantly more}

precipitation in north-western India (up to 0.25 mm day−1_{). Most striking is that}

pre-cipitation is significantly increased by up to 0.45 mm day−1 _{in the Shaanxi-Shandong}

region, where the prehistoric cultures settled.

To better visualise the different effect of afforestation under 6k- and 0k-orbital

con-5

ditions, Fig. 14b illustrates the di_fference of the a_fforestation-induced precipitation anomalies (6k–0k). It emphasises that afforestation leads to more precipitation in south- and central-eastern China as well as in India when mid-Holocene instead of present-day orbit is prescribed. These di_fferences in precipitation signal can be related to a change in circulation and moisture convergence (Fig. 14d). With prescribed

6k-10

orbit, afforestation leads to a stronger monsoon flow anomaly onto the continent than under 0k-orbital conditions. Therefore, more moisture is transported from the Pacific Ocean and Bay of Bengal to East China and India.

Figure 14c shows the monsoon precipitation difference associated with the pure or-bital forcing. Compared to this signal, the effect of afforestation on monsoon

precipita-15

tion is weak. However, in the region with prehistoric settlement in eastern China, the precipitation change due to afforestation (6k) is approx. half as large as the change induced by the orbital forcing. Averaged over the region of ca. 99◦_E–120◦_{E and 28}◦_N–

42◦

N, afforestation leads to a precipitation increase of 0.14 mm day−1

. The orbital forcing results in 0.25 mm day−1_{additional precipitation.}

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According to our model results, large-scale Holocene forest decline, regardless, whether it has been climate- or human-induced, could thus strongly have contributed to the decrease of summer monsoon precipitation reconstructed for the East Asian monsoon region (e.g. Wang et al., 2005). However, the magnitude of the assumed for-est loss in our study is intentionally exaggerated. The resulting climate signal can

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additional environmental pressure on the Asian human civilisations by further reducing water availability.

5 Summary and conclusion

This study highlights the effect large-scale land cover change in the Asian monsoon domain could have on the regional and remote climate. By applying the general

5

circulation model ECHAM5/JSBACH a set of numerical experiments has been performed with either complete forest cover or complete grass cover or potential present-day vegetation prescribed. Our analyses address the following questions: How do afforestation and deforestation influence the local climate? How large is the remote effect on precipitation in North Africa and the Middle East? Could early- and

10

mid-Holocene land cover changes have contributed to the decreasing Asian monsoon precipitation known from reconstructions?

a. Local effect:

– The impact of land cover change on local precipitation is most pronounced in the

15

monsoon season. During this time, deforestation leads to a broad reduction of precipitation in the Asian monsoon domain. In contrast, the change of precip-itation resulting from afforestation is arranged in an alternating pattern: rainfall is decreased in South India and the Great Plain of China, while it is increased in south-eastern China, the southern part of the Tibetan Plateau and Northeast

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China.

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– The impact of land cover change on local near-surface air temperatures is most pronounced in the dry/cold season. During this time, deforestation leads to a cooling and afforestation to a warming of the climate in most parts of the Asian monsoon domain. In the monsoon season, the signal is of similar shape in the regions with a rather temperate or cool climate, i.e. the northern parts. In the

5

southern parts where climate is tropical, afforestation (deforestation) leads to a temperature decrease (increase) resulting from an enhanced (reduced) evapora-tive cooling.

– Most of the temperature signal can be explained directly by the change in sur-face albedo due to the land cover change and the accompanied alteration of the

10

surface energy balance. b. Remote effect:

– Changes in the Asian monsoon system can affect the climate in North Africa and the Middle East via the modification of the large-scale circulation. In our experiments, the remote effect of land cover change in the Asian monsoon domain

15

on the precipitation pattern in these regions is strong with respect to the relative precipitation amount.

– According to our simulations, deforestation in the Asian monsoon domain signifi-cantly increases precipitation in most parts of North Africa. In some parts of the Sahara, rainfall is even doubled.

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c. Role of Holocene land-cover change:

– The amplitude of the climate signal induced by afforestation does not change when prescribing mid-Holocene instead of present-day orbital conditions, but the pattern shifts, showing more precipitation in the region along the Yellow River. – Compared to the climate change related to the orbital forcing, the overall climate

5

signal induced by the land cover change is small in our model experiments. How-ever, in the settlement area of East Asian prehistoric societies (mainly along the Yellow River) the precipitation change associated with the a_fforestation is half as large as the orbitally-induced precipitation change. In other words: under the as-sumption that the entire Asian monsoon region was covered by forest in

early-10

and mid-Holocene, the loss of forest to the present-day potential land cover is responsible for more than one-third of the total mid- to late Holocene precipitation change in the settlement area of the major East Asian neolithic civilisations. – The here prescribed complete forest cover is definitely an exaggerated

represen-tation of the mid-Holocene land cover in the Asian monsoon region. Therefore,

15

the amplitude of the climate signal is probably too high.

– Our model results show that large-scale forest decline in East and South Asia leads to heavy losses in regional precipitation. One can not exclude that climate-or human-induced land cover changes have contributed to the decline of majclimate-or prehistoric civilisations by amplifying the general orbital-induced attenuating Asian

20

summer monsoon.

Acknowledgements. This research was funded by the German Research Foundation (DFG, grant CL178/6-1,2). The authors would like to thank Victor Brovkin, MPI-M, for constructive discussion.

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References

Ansari, M. and Vink, A.: Vegetation history and palaeoclimate of the past 30 kyr in Pakistan as inferred from the palynology of continental margin sediments off the Indus Delta, Rev. Paleobot. Palynol., 145, 201–216, 2007.

Bathiany, S., Claussen, M., Brovkin, V., Raddatz, T., and Gayler, V.: Combined biogeophysical

5

and biogeochemical effects of large-scale forest cover changes in the MPI earth system model, Biogeosciences, 7, 1383–1399, doi:10.5194/bg-7-1383-2010, 2010.

Bonan, G. B.: Forests and climate change: Forcings, feedbacks, and the climate benefits of forests, Science, 320, 1444-1449, 2008.

Bonan, G. B., Pollard, D., and Thompson, S. L.: Effects of boreal forest vegetation on global

10

climate, Nature, 359, 716-718, 1992.

Chalita, S. and Le Treut, H.: The albedo of temperate and boreal forest and the Northern Hemisphere climate: a sensitivity experiment using the LMD GCM, Climate Dyn., 10, 231-240, 1994.

Claussen, M., Brovkin, V., and Ganopolski, A.: Biogeophysical versus biogeochemical

feed-15

backs of large scale land cover change, Geophys. Res. Lett., 28, 1011–1014, 2001.

Clift, P. D. and Plumb, R. A.: The Asian monsoon: causes, history and e_ffects, Cambridge Univ. Press, Cambridge, 270 pp., 2008.

Cui, X., Graf, H. F., Langmann, B., Chen, W., and Huang, R.: Climate impacts of anthropogenic land use changes on the Tibetan Plateau, Global Planet Change, 54, 33–56, 2006.

20

Dallmeyer, A., Claussen, M., and Otto, J.: Contribution of oceanic and vegetation feedbacks to Holocene climate change in monsoonal Asia, Clim. Past, 6, 195–218, doi:10.5194/cp-6-195-2010, 2010.

Duan, A. and Wu, G. X.: Role of the Tibetan Plateau thermal forcing in the summer climate pattern over subtropical Asia, Climate Dyn., 24, 793–807, 2005.

25

Findell, K. L., Knutson, T. R., and Milly, P. C. D.: Weak simulated extratropical responses to complete tropical deforestation, J. Climate, 19, 2835–2850, 2006.

Fleitmann, D., Burns, S. J., Mudelsee, M., Neff, U., Kramers, J., Mangini, A., and Matter, A.: Holocene forcing of the Indian monsoon recorded in a stalagmite from Southern Oman, Science, 300, 1737–1739, 2003.

30

(23)

BGD

8, 1693–1731, 2011

Title Page

Tables Figures

◭ ◮

Back Close

Full Screen / Esc

Discussion

P

a

per

|

Dis

cussion

P

a

per

|

Discussion

P

a

per

|

Discussio

n

P

a

per

|

Gao, X. J., Zhang, D. F., Chen, Z. X., Pal, J. S., and Giorgi, F.: Land use effects on climate in China as simulated by a regional climate model, Science in China Series D, Earth Sciences, 50(4), 620–628, 2007.

Gedney, N. and Valdes, P. J.: The effect of Amazonian deforestation on the Northern Hemi-sphere circulation and climate, Geophys. Res. Lett., 27, 3053–3056, 2000.

5

Hasler, N., Werth, D., and Avissar, R.: Effects of tropical deforestation on global hydroclimate: A multimodel ensemble analysis, J. Climate, 22, 1124–1141, 2009.

Henderson-Sellers, A., Dickinson, R. E. , Durbidge, T. B., Kennedy, P. J., McGuffie, K., and Pitman, A. J.: Tropical deforestation: Modeling local-scale to regional-scale climate change, J. Geophys. Res., 98(D4), 7289–7315, 1993.

10

Herzschuh, U., Birks, H. J. B., Ni, J., Zhao, Y., Liu, H., Liu, X., and Grosse, G.: Holocene land-cover changes on the Tibetan Plateau, Holocene, 20, 91–104, 2010.

Ivory, S. and L ´ezine, A.-M.: Climate Environmental change at the end of the Holocene Humid Period: a pollen record o_ffPakistan, C. R. Geosci., 341, 760–769, 2009.

Jungclaus, J. H., Botzet, M., Haak, H., Keenlyside, N., Luo, J.-J., Latif, M., Marotzke, J.,

Mikola-15

jewicz, U., and Roeckner, E.: Ocean circulation and tropical variability in the coupled model ECHAM5/MPI-OM, J. Climate, 19, 3952–3972, 2006.

Li, C. and Yanai, M.: The onset and interannual variability of the Asian summer monsoon in relation to land-sea thermal contrast, J. Clim., 9, 358–375, 1996.

Maher, B. A.: Holocene variability of the East Asian summer monsoon from Chinese cave

20

records: a re-assessment, Holocene, 18(6), 861–866, 2008.

Marsland, S. J., Haak, H., Jungclaus, J. H., Latif, M., and Roske, F.: The Max-Planck-Institute global ocean/sea ice model with orthogonal curvilinear coordinates, Ocean Modell., 5, 91– 127, 2003.

Pielke, R. A., Avissar, R., Raupach, M., Dolman, A. J., Zeng, X. B., and Denning, A. S.:

Interac-25

tions between the atmosphere and terrestrial ecosystems: influence on weather and climate, Glob. Change Biol., 4, 461–475, 1998.

Polcher, J. and Laval, K.: The impact of African and Amazonian deforestation on tropical cli-mate, J. Hydrol., 155, 389–405, 1994.

Pongratz, J., Reick, C., Raddatz, T., and Claussen, M.: A reconstruction of global

agricul-30

tural areas and land cover for the last millennium, Global Biogeochem. Cyc., 22, GB3018, doi:10.1029/2007GB003153, 2008.

(24)

BGD

8, 1693–1731, 2011

Title Page

Tables Figures

◭ ◮

Back Close

Full Screen / Esc

Discussion

P

a

per

|

Dis

cussion

P

a

per

|

Discussion

P

a

per

|

Discussio

n

P

a

per

|

Wetzel, P., and Jungclaus, J.: Will the tropical land biosphere dominate the climate-carbon cycle feedback during the twenty-first century?, Clim. Dyn., 29, 565–574, 2007.

Ren, G.: Decline of the mid- to late Holocene forests in China: climatic change or human impact?, J. Quat. Sci., 15(3), 273–281, 2000.

Ren, G.: Changes in forest cover in China during the Holocene, Veget. Hist. Archaeobot., 16,

5

119–126, 2007.

Ren, G. and Beug, H-J.: Mapping Holocene pollen data and vegetation of China, Quat. Sci. Rev., 21, 1395–1422, 2002.

Rodwell, M. J. and Hoskins, B. J.: Monsoons and the dynamics of deserts, Quart. J. Roy. Meteor. Soc., 122, 1385–1404, 1996.

10

Roeckner, E., B ¨auml, G., Bonaventura, L., Brokopf, R., Esch, M., Giorgetta, M., Hagemann, S., Kirchner, I., Kornblueh, L., Manzini, E., Rhodin, A., Schlese, U., Schultzweida, U., and Tompkins, A.: The atmospheric general circulation model ECHAM5, Part I: Model descrip-tion, MaxPlanckInst. f. Meteor., Report No. 349, Hamburg, 2003.

Singh, R., Wasson, R. J., and Agrawal, D. P.: Vegetational and seasonal climate changes

15

since the last full glacial in the Thar Desert, northwestern India, Rev. Paleobot. Palynol., 64, 351–358, 1990.

Snyder, P. K.: The influence of tropical deforestation on the Northern Hemisphere climate by atmospheric teleconnections, Earth Interact., 14, 1–34, doi:10.1175/2010EI280.1, 2010. Wang, Y. J., Cheng, H., Edwards, R. L., He, Y. Q., Kong, X. G., An, Z. S., Wu, J. Y., Kelly, M.

20

J., Dykoski, C. A., and Li, X. D.: The Holocene Asian monsoon: Links to solar changes and North Atlantic climate, Science, 308, 854–857, 2005.

Webster, P. J., Maga `oa, V. O., Palmer, T. N., Shukla, J., Tomas, R. A., Yanai, M., and Yasunari, T.: Monsoons: Processes, predictability, and the prospects for prediction, J. Geophys. Res., 103(C7), 14451–14510, 1998.

25

Werth, D. and Avissar, R.: The local and global effects of Amazonian deforestation, J. Geophys. Res., 107, 8087, doi:10.1029/2001JD000717, 2002.

Werth, D. and Avissar, R.: The local and global effects of Southeast Asian deforestation, Geo-phys. Res. Lett., 32, L20702, doi:10.1029/2005GL022970, 2005a.

Werth, D. and Avissar, R.: The local and global effects of African deforestation, Geophys. Res.

30

Lett., 32, L12704, doi:10.1029/2005GL022969, 2005b.

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Tables Figures

◭ ◮

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Full Screen / Esc

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P

a

per

|

Dis

cussion

P

a

per

|

Discussion

P

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n

P

a

per

|

Ye, D. Z. and Wu, G. X.: The role of the heat source of the Tibetan Plateau in the general circulation, Meteorol. Atmos. Phys., 67, 181–198, 1998.

Yu, G., Chen, X., Ni, J., Cheddadi, R., Guiot, J., Han, H., Harrison, S. P., Huang, C., Ke, M., Kong, Z., Li, S., Li, W., Liew, P., Liu, G., Liu, J., Liu, Q., Liu, K.-B., Prentice, I. C., Qui, W., Ren, G., Song, C., Sugita, S., Sun, X., Tang, L., Van Campo, E., Xia, Y., Xu, Q., Yan, S.,

5

Yang, X., Zhao, J., and Zheng, Z.: Palaeovegetation of China: a pollen date-based synthesis for the mid-Holocene and last glacial maximum, J. Biogeography, 27, 635–664, 2000. Zhao, Y., Yu, Z., and Chen, F.: Spatial and temporal patterns of Holocene vegetation and

climate changes in arid and semi-arid China, Quat. Int., 194, 6–18, 2009.

Zhang, H. Q. and Gao X. J.: On the atmospheric dynamical response to land-use change in

10

East Asian monsoon region, Clim. Dyn., 33, 409–426, 2008.

Zheng, Y. Q., Yu, G., Qian, Y. F., Miao, M., Zeng, X. M., and Liu, H. Q.: Simulations of regional climatic e_ffects of vegetation change in China, Q. J. R. Meteorol. Soc., 128, 2089–2114, 2002.

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Fig. 3. Simulated precipitation change between the afforestation (for) or deforestation (gra) experiment and the control run (ctrl), averaged over the monsoon season (left panel) and the dry/cold season (right panel). The monsoon season lasts from May to October, the dry/cold season contains the month from November to April. Values are given in mm day−1_{. Only values}

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Fig. 5.Simulated change of different parameters influencing the precipitation change between the afforestation experiment (for) and the control run (ctrl), averaged over the monsoon season.

(a) total cloud cover in %; (b) evaporation in mm day−1_; _(c) _{mean sea level pressure in Pa,}

shaded, combined with 850 hPa wind in m s−1_{, vector;}_(d)_{precipitation-evaporation in mm day}−1

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Fig. 9.Simulated change of different parameters influencing the temperature change between the a_fforestation experiment (for) and the control run (ctrl), averaged over the monsoon season.

(a)albedo in %; (b)net surface shortwave radiation in W m−2_; _(c)_{latent heat flux in W m}−2_;

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